Lithos 208–209 (2014) 53–66

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Geochemistry of ultrapotassic volcanic rocks in Xiaogulihe NE China:Implications for the role of ancient subducted sediments

Yang Sun a,b, Jifeng Ying a,⁎, Xinhua Zhou a, Ji'an Shao c, Zhuyin Chu a, Benxun Su a

a State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, Chinab College of Earth Science, University of Chinese Academy of Sciences, Beijing 100049, Chinac Key Laboratory of Orogenic Belts and Crustal Evolution, School of Earth and Space Sciences, Peking University, Beijing 100871, China

⁎ Corresponding author. Tel.: +86 10 82998532; fax: +E-mail address: jfying@mail.iggcas.ac.cn (J. Ying).

http://dx.doi.org/10.1016/j.lithos.2014.08.0260024-4937/© 2014 Elsevier B.V. All rights reserved.

a b s t r a c t

a r t i c l e i n f o

Article history:Received 19 May 2014Accepted 29 August 2014Available online 16 September 2014

Keywords:Ultrapotassic igneous rocksSr–Nd–Hf–Pb–Os isotopesSub-continental lithospheric mantleAncient subducted sedimentsNE China

The unique eruptions of ultrapotassic volcanic rocks in eastern China reported so far took place in theXiaogulihe area of western Heilongjiang Province, NE China. These ultrapotassic rocks are characterizedby extremely high K2O contents (N7 wt.%), abnormally unradiogenic Pb isotopic compositions (206Pb/204Pb =16.44–16.55; 207Pb/204Pb = 15.39–15.46; 208Pb/204Pb = 36.35–36.61), and moderately high 87Sr/86Sr ratios(0.7053–0.7057), which can be basically correlated with those of ultrapotassic igneous rocks distributed widelyin northwestern America and Aldan Shield. The positive correlation between 187Os/188Os and 1/Os argues thatthese ultrapotassic rocks have probably experienced negligible lower continental crust addition (less than 1%)during magma ascent. The high contents of K2O and negative correlation between 87Sr/86Sr and 206Pb/204Pb ofthese ultrapotassic rocks indicate the presence of a potassic phase, mostly phlogopite, in their mantle source.Their strong fractionation of rare earth elements and lack of Nd–Hf isotopic decoupling reveal a low-degreepartial melting of garnet-bearing source rocks. In addition, the low CaO and Al2O3 contents of whole-rockcompositions and low Fe/Mn ratios of olivine phenocryst chemistries suggest peridotites rather than pyroxenitesas dominant source rocks for the Xiaogulihe ultrapotassic rocks. Based on these distinctive geochemical charac-teristics, we thus propose that the ultimate mantle source of the Xiaogulihe ultrapotassic volcanic rocks isphlogopite-bearing garnet peridotite within the lower part of the sub-continental lithospheric mantle(SCLM) that had been metasomatized by potassium-rich silicate melts. Combined with the unradiogenic Pbcompositions, the most likely source of these potassium-rich silicate melts is the ancient subductedcontinental-derived sediments (N1.5 Ga). These ancient subducted sediments, possessing relatively low initialPb isotopic compositions, had experienced large U/Pb fractionation during a subduction process, resulting inlow-μ (238U/204Pb), and then accumulated in the mantle transition zone. The relatively low 87Sr/86Sr ratios ofthese ultrapotassic rocks also imply that their mantle source had evolved with low Rb/Sr ratios, which possiblyresulted from the metasomatized melts derived from the ancient subducted sediments. This interpretation isquite different from previous hypotheses that attribute their unusual geochemical features to a dominantlyasthenospheric source with a contribution from delaminated ancient SCLM, or a SCLM source that has beenmetasomatized by melts derived from deep asthenosphere or delaminated ancient lower continental crust.

© 2014 Elsevier B.V. All rights reserved.

1. Introduction

Interest in the petrogenesis of ultrapotassic igneous rocks increasedgreatly in the late 1970s with the discovery of diamond-bearing olivinelamproites in NWAustralia (Atkinson et al., 1984). Because of their un-usual geochemistry, distinctive mineralogy and potential to constrainthe origin of an enriched mantle (EM) component, numerous workshave been done to explain the magma genesis and mantle sourceevolution of these ultrapotassic rocks over the past 30 years(Avanzinelli et al., 2008; Chen et al., 2007; Chu et al., 2013; Davies

86 10 62010846.

et al., 2006; Foley and Peccerillo, 1992; Foley et al., 1987; Fraser et al.,1985; Kuritani et al., 2013; Mitchell and Bergman, 1991; Murphyet al., 2002; Nelson et al., 1986; Prelevic et al., 2008; Zhang et al.,1995; Zou et al., 2003).

Considering the distinctive geochemical characteristics of theseultrapotassic rocks, none of the previous studies have successfully andconvincingly explained the nature and evolutionary history of theirmantle sources, and many questions remain to be solved, such as thesite of mantle sources, i.e. whether their ultimate sources are locatedwithin the SCLM (Avanzinelli et al., 2008; Chen et al., 2007; Chu et al.,2013; Davies et al., 2006; Prelevic et al., 2008; Zhang et al., 1995; Zouet al., 2003), asthenosphere (Choi et al., 2006) or the mantle transitionzone (Kuritani et al., 2013; Murphy et al., 2002); and where the fluids

54 Y. Sun et al. / Lithos 208–209 (2014) 53–66

and/or melts which metasomatized the ultimate mantle sources camefrom, i.e. deep asthenospheric mantle (McKenzie, 1989; Zhang et al.,2000), subduction zone (Elburg and Foden, 1999), delaminated SCLM(Choi et al., 2006; Zhao et al., 2014), mantle transition zone (Kuritaniet al., 2013; Murphy et al., 2002) or delaminated lower continentalcrust (Chu et al., 2013); and also the source lithology of thepotassium-rich rocks, i.e. whether they are peridotite or pyroxenite(Foley, 1992).

Cenozoic volcanic rocks, mostly alkaline basalts, are widely distrib-uted inNE China and forman important part of theWest Pacific volcaniczone (Basu et al., 1991; Chen et al., 2007; Choi et al., 2006; Chu et al.,2013; Kuritani et al., 2013; Liu et al., 1994; Zhang et al., 1995; Zhouand Armstrong, 1982; Zou et al., 2003). Among them, there are threepotassic volcanic areas of Wudalianchi (WDLC), Erkeshan (EKS) andKeluo, called the WEK potassic volcanic field (Zhang et al., 1991).Detailed geochemical studies have revealed that the WEK potassicrocks displayed a clear EMI-like signature (Chen et al., 2007; Chuet al., 2013; Kuritani et al., 2013; Zhang, 1992; Zhang et al., 1991,1995, 1998; Zou et al., 2003). Moreover, there is another volcanicfield, around 200 km northwest to the WEK potassic volcanic field,called Xiaogulihe (Shao et al., 2009). The rocks here, having extremelyhigh K2O contents (over 7 wt.%) and K2O/Na2O ratios (over 2), can beclassified as ultrapotassic rocks following the definition of Foley et al.(1987). Few investigations have been carried out on these ultrapotassicrocks so far, although they are distinct from all the Cenozoic basalts ineastern China in terms of their extremely high K2O contents, mostlyenriched incompatible elements, and very typical EMI-like isotopicsignatures (especially the least radiogenic Pb isotopic compositions).In the following discussion, the term ‘WEK potassium-rich volcanicrocks’will be used to substitute for the Xiaogulihe ultrapotassic volcanicrocks and the WEK potassic volcanic rocks.

In this study, we present, for the first time, major-, trace- andplatinum group element (PGE) abundances and Sr–Nd–Hf–Pb–Os sys-tematic isotopic compositions of the Xiaogulihe ultrapotassic volcanicrocks with aims to further understand the range of chemical andgeodynamic processes that have contributed to the petrogenesis of

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120 E 125 E 130 E 135 E 140 E 145 E

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Fig. 1. (a) Tectonic framework of northeast China and Far East Russia (modified after Zhou anddanjiang Fault; F4 = Heilongjiang Fault; F5 = Hegenshan–Heihe Fault; F6 = Xinlin–Xiguitu Fasketch map showing major faults and the location of the Xiaogulihe, Keluo, Wudalianchi and E

these rocks. Consistent with previous studies, our results favor that theXiaogulihe ultrapotassic rocks mainly originated from the lower SCLMthat has been enriched through an ancient or recent metasomaticevent. Moreover, we propose that the metasomatic melts were mainlyderived from ancient continental-derived sediments subducted intothe mantle transition zone with an oceanic plate. The metasomatizedSCLM then experienced low-degree partial melting following a recentupwelling of asthenosphere and crustal attenuation in the SongliaoBasin or continental rifting in theWEK potassium-rich volcanic rock belt.

2. Geological setting and sample descriptions

The WEK potassium-rich volcanic field, including the Xiaoguliheultrapotassic volcanic field and the WEK potassic volcanic field, islocated at the boundary between the northwestern margin of theSongliao Basin and the Great Xing'an Ranges, both of which are withinthe Xing'an–Mongolia Orogenic Belt (XMOB) (Fig. 1a). This NNW-trending, 400 km-long Cenozoic volcanic rock belt is one of the mainpotassium-rich volcanism areas in China.

The XMOB (Fig. 1a), generally considered as the eastern segment ofthe Central Asian Orogenic Belt, links the Siberian Craton in the northand the Sino-Korean Craton in the south (Li, 2006; Sengor and Natalin,1996; Sengor et al., 1993; Tang et al., 2014; Xu et al., 2013; Zhou andWilde, 2013). It is surrounded by the Paleo-Pacific orogens includingthe Mongolian–Okhotsk orogens in the north and Russian Far Eastorogens in the east, and separated from the Sino-Korean Cratonby the Solonker–Xar Moron–Changchun–Yanji fault belt in the South(Li, 2006; Xu et al., 2013). The tectonic evolution of the XMOB wascharacterized by several significant events, including the evolutionand final closure of the Paleo-Asian Ocean, the amalgamation of severalmicro-continental massifs (e.g. from west to east, the Erguna, Xing'an,Songliao, Jiamusi–Khanka massifs), and the subduction of the Paleo-Pacific Ocean (Li, 2006; Sengor and Natalin, 1996; Sengor et al., 1993;Tang et al., 2014; Xu et al., 2013).

The upwelling of asthenosphere and attenuation of continental crustin the Songliao Basin have been demonstrated by seismic, electric

125 E 130 E

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Cenozoic basalts

Potassium-richvolcanic rocks

Major faultLesser Xing’an Mts

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China

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Wilde, 2013). F1 = Solonker–Xar Moron–Changchun zone; F2 = Yanji Fault; F3 = Mu-ult; F7= Yilan–Yitong Fault; F8= Dunhua–Mishan Fault, and F9= Primoria Fault. (b) Arkeshan potassium-rich volcanic rocks in NE China (modified after Zhang et al., 2000).

55Y. Sun et al. / Lithos 208–209 (2014) 53–66

conductivity and gravity measurements. The lithospheric thickness atthe center of Songliao Basin is about 80 km, 40–70 km less than thatat the margins, and the minimum crustal thickness beneath the basinis 23 km, about 10–17 km thinner than that of the surrounding areas(Hsü, 1989; Ma, 1987; Yuan, 1996). Both earthquake focal mechanismsolutions (Ma, 1987) and inversion of geopotential data (Liu, 1978)suggest that the WEK potassium-rich volcanic field is under tensilestress at present.

Fresh sampleswere collected from lava flows related to the eruptionof Ma'anshan (K–Ar, 0.19–0.30 Ma, Zhang, 1992) from the Xiaoguliheultrapotassic volcanic field. Olivine, clinopyroxene, and leucite pheno-crysts are distributed in the groundmass, while the matrix is consistedof olivine, clinopyroxene, leucite, nepheline, sodalite, feldspar, phlogo-pite, and Fe–Ti oxide. The microphotographs of representative samplesare shown in Fig. 2.

3. Analytical methods

All samples selected for whole-rock major, trace elemental andisotopic analyses were first trimmed into small chips to completelyremove the weathered surface. Then, the samples were further splitinto smaller chips by a hammer wrapped in soft cloths. Only smallchips that were completely free of surface alteration and mantleor/and crustal xenocrysts were selected and cleaned with hydrochloricacid and deionized water (see Chu et al., 2013 for details). The cleanedrock chipswere then crushedwith a ceramic jaw crusher and powderedusing an agate mill to ~200 mesh.

3.1. Major and trace elemental analyses

Whole-rock major elemental compositions were analyzed byX-ray fluorescence spectrometry (XRF) using an Axios-Minerals spec-trometer at the Institute of Geology and Geophysics, Chinese Academyof Sciences (IGGCAS), following the procedures of Chu et al. (2009).Fused glass disks were used and the precisions are 1–3% RSD forelements' concentrations N 1 wt.%, and about 10% RSD for elements'concentrations b 1 wt.%. Loss on ignition (LOI) was determined byroutine procedures. Chinese National Standard GSR-3 was measuredunder the same analytical conditions and the values are in goodagreement with reference values within the analytical errors (Table 1).

Trace elemental abundances were measured by inductively coupledplasma mass spectrometry (ICP-MS) at IGGCAS using a Finnigan MATElement spectrometer, following the procedures described in Chuet al. (2009). GSR-3was analyzed to evaluate the precision and accuracyof the analytical procedures, and the results are well consistent withrecommended values within the analytical errors (Table 1). Both preci-sion and accuracy are generally better than 5% for most trace elements.

Fig. 2. Representative microphotographs of Xiaogulihe ultrapotassic volcanic rocks. (a) Samplecryst aggregate and leucite phenocrysts.

3.2. Sr–Nd–Hf–Pb isotopic analyses

The Rb–Sr and Sm–Nd isotopic compositions were determined atIGGCAS using a GV Isoprobe-T thermal ionization mass spectrometer,following the procedures described in Chu et al. (2009). The pro-cedure involved Rb, Sr and REEs separation with AG 50 W × 12(200–400 mesh) columns, and Eichrom®-LN (50–100 m) columnsfor separating Nd and Sm from other REEs. Measured 87Sr/86Sr and143Nd/144Nd ratios were corrected for mass fractionation by normaliz-ing to 86Sr/88Sr = 0.1194 and 146Nd/144Nd = 0.7219, respectively.During the period of data collection, the measured values for the NBS-987 Sr standard and the JNdi-1 Nd standard were 87Sr/86Sr =0.710246 ± 16 (2SD, n = 8) and 143Nd/144Nd = 0.512118 ± 11 (2SD,n = 8), respectively. The Lu–Hf isotopic compositions were conductedby a Thermo Fisher Scientific NEPTUNEMC-ICP-MS at IGGCAS, followingthe procedures described in Yang et al. (2010). The Eichrom®-LN(100–150m) columnswere used for Lu–Hf separation. Hf isotopic ratioswere normalized to 179Hf/177Hf= 0.7325 and 176Lu/175Lu isotopic ratioswere normalized by Yb isotopic ratios. During the analysis procedure,the measured values for the Alfa Hf standard were 176Hf/177Hf =0.282185 ± 5 (2SD, n = 6). The USGS basalt standards BCR-2 andBHVO-2 were analyzed for Sr–Nd–Hf isotopic compositions to evaluatethe reliability of the analytical procedures, and the results correspondwell with the reported literature values (Table 2). All the chemicalprocedures were carried out in a clean laboratory environment;procedural blanks were about 40 pg for Rb, 300 pg for Sr, 20 pg for Sm,60 pg for Nd, 10 pg for Lu and 40 pg for Hf.

The Pb isotopic compositions were carried out at IGGCAS using aFinnigan MAT262 thermal ionization mass spectrometer. The whole-rock powderswere dissolved in purifiedHF+HNO3 and then separatedusing anion-exchange columns with diluted HBr as elutant. During theanalytical session, repeated analyses for the NBS-981 Pb standardyielded 206Pb/204Pb = 16.942 ± 14 (2SD, n = 6), 207Pb/204Pb =15.497 ± 14 (2SD, n = 6), 208Pb/204Pb = 36.729 ± 14 (2SD, n = 6).

3.3. Re–Os and PGE analyses

The Re–Os isotopic compositions and PGE abundances were mea-sured by isotope dilution methods at IGGCAS, following the proceduresdescribed in Chu et al. (2013). Simply, about 2 g sample powder togeth-er with suitable amounts of mixed Re–Os spike (187Re–190Os) and PGEspike (191Ir–99Ru–194Pt–105Pd), and inverse aqua regia (3 ml 12 N HCland 6 ml 15 N HNO3) were added to the Carius tube, and digested atabout 250 °C for 48–72 h in an oven. Osmium was first extracted fromthe aqua regia solution by CCl4 and then back-extracted byHBr, and fur-ther purified throughmicro-distillation following themethod describedin Birck et al. (1997). The total procedural Os blank was 0.1–0.3 pgwith 187Os/188Os of about 0.16, which was negligible for all samples.

GLH12-02, euhedral leucite phenocrysts; (b) sample GLH12-16, euhedral olivine pheno-

Table 1Elemental compositions of the Xiaogulihe ultrapotassic volcanic rocks.

Sample GLH12-02 GLH12-03 GLH12-04 GLH12-05 GLH12-06 GLH12-07 GLH12-08 GLH12-09 GLH12-10 GLH12-11 GLH12-11R GLH12-12

Major element (wt.%)SiO2 49.31 48.91 49.02 49.65 48.76 48.12 48.64 48.45 47.93 47.98 47.91 48.73TiO2 3.20 3.17 3.18 3.18 3.08 3.00 3.12 3.05 3.03 3.03 3.02 3.04Al2O3 11.06 10.98 10.95 11.19 10.88 11.11 11.22 10.97 11.27 11.01 11.03 10.87TFe2O3

b 7.98 7.93 8.18 8.18 8.37 8.67 8.50 8.35 8.66 8.64 8.61 8.29MnO 0.11 0.11 0.11 0.11 0.11 0.11 0.13 0.13 0.12 0.11 0.12 0.11MgO 7.74 7.83 7.84 8.05 8.74 8.88 8.35 8.64 8.82 9.02 8.99 8.67CaO 5.65 5.72 6.19 5.78 6.18 6.56 6.34 6.49 6.67 6.66 6.65 6.45Na2O 2.78 2.60 2.45 2.06 2.39 1.71 1.88 2.09 1.96 2.60 2.60 2.42K2O 9.29 9.26 8.86 9.08 8.37 7.82 8.42 8.01 7.81 7.80 7.78 8.58P2O5 1.83 1.85 1.81 1.35 1.48 1.73 1.89 1.85 1.79 1.77 1.77 1.30LOI 0.26 0.26 0.94 0.70 1.02 1.28 0.94 1.34 1.04 0.76 0.74 1.06Total 99.21 98.62 99.53 99.33 99.38 98.99 99.43 99.37 99.10 99.38 99.22 99.52K2O/Na2O 3.34 3.56 3.62 4.41 3.50 4.57 4.48 3.83 3.98 3.00 2.99 3.55Mg# 69.3 69.7 69.1 69.6 70.9 70.5 69.6 70.7 70.4 70.9 70.9 70.9Or 29.14 29 40 43.57 37.64 47.35 50.6 48.39 47.14 34.59 34.95 33.46Ne 2.84 2.72 3.85 3.85 5.17 3.7 5.02 6.11 5.89 7.28 7.4 4.49Lc 20.82 21.05 10.33 8.57 10 0 0 0 0 9.64 9.31 14.2Di 12.75 13.03 10.18 11.8 14.18 12.01 11.23 13.75 13.32 17 16.98 14.06Ol 10.2 10.63 10.59 10.43 10.94 11.98 11.17 10.95 11.4 10.49 10.46 10.82Ac 12.86 12.22 12.29 9.33 9.75 0.78 4.55 5.27 1.89 7.88 7.7 11.05Il 6.16 6.14 2.8 3.41 4.23 2.38 3.3 4.34 3.29 5.85 5.84 2.69Ap 4.3 4.37 4.26 3.18 3.49 4.11 4.45 4.38 4.24 4.17 4.18 3.06

Trace element (ppm)Sc 20.1 19.5 18.9 18.6 18.7 18.5 18.3 17.9 18.5 18.0 17.7 18.8Cr 268 271 274 239 265 336 286 282 313 336 331 300Ni 174 178 181 183 200 193 186 197 199 194 192 200Cu 26.6 31.8 36.5 38.9 39.8 32.5 42.3 34.4 39.1 39.4 39.5 36.6Zn 111 114 116 115 114 109 119 116 107 114 111 111Rb 162 163 156 156 146 142 151 191 204 135 134 152Sr 1827 1828 1513 1002 1233 1244 1598 1406 1732 1756 1762 1194Y 23.7 23.8 24.4 24.3 24.2 22.3 26.3 24.1 23.7 22.8 22.9 23.3Zr 871 869 931 905 876 748 886 842 747 743 733 829Nb 64.2 62.0 65.8 65.7 66.5 64.4 65.4 63.6 65.0 64.3 63.2 65.6Ba 3599 3615 3512 2131 3029 2437 3265 3328 3706 3301 3320 3551La 163 167 168 172 161 146 167 158 148 146 146 153Ce 312 316 320 316 308 276 313 298 278 271 275 296Pr 35.6 36.1 36.7 37.5 35.7 31.6 35.7 34.3 32.4 31.1 31.4 33.9Nd 120 124 125 128 124 108 122 116 111 106 107 119Sm 18.6 19.1 19.7 19.8 18.7 17.2 19.1 18.4 17.4 16.9 17.0 17.8Eu 4.04 4.23 4.32 4.32 4.19 3.88 4.32 4.03 3.89 3.78 3.82 4.00Gd 12.5 13.3 13.4 13.5 13.2 12.1 13.8 12.8 12.3 11.9 12.0 12.6Tb 1.36 1.45 1.48 1.48 1.46 1.35 1.50 1.39 1.36 1.32 1.34 1.40Dy 5.80 6.03 6.41 6.13 6.30 6.00 6.31 6.17 6.00 5.89 5.94 6.15Ho 0.947 0.957 1.03 1.03 1.03 0.963 1.02 0.983 0.978 0.952 0.958 0.982Er 2.16 2.14 2.31 2.37 2.41 2.22 2.31 2.23 2.27 2.20 2.23 2.28Tm 0.261 0.276 0.295 0.298 0.291 0.275 0.296 0.279 0.279 0.274 0.269 0.275Yb 1.49 1.61 1.68 1.64 1.69 1.58 1.73 1.51 1.50 1.59 1.55 1.48Lu 0.206 0.217 0.228 0.222 0.224 0.215 0.239 0.204 0.209 0.216 0.213 0.204Hf 20.9 22.6 24.4 23.9 22.9 19.7 23.0 22.1 19.6 19.5 19.3 21.9Ta 3.20 3.34 3.59 3.66 3.65 3.59 3.55 3.45 3.63 3.58 3.57 3.59Pb 28.9 31.3 31.0 29.6 10.6 21.0 32.6 18.9 25.0 30.2 30.2 24.3Th 7.71 8.34 9.07 8.83 9.11 8.74 9.15 8.75 8.55 8.58 8.54 8.88U 1.42 1.59 1.67 1.47 1.64 1.59 1.67 1.55 1.49 1.58 1.56 1.44∑REE 679.1 692.3 700.2 704.1 678.2 606.7 688.3 654.2 615.0 599.6 605.1 648.2

Element ratiosCe/Pb 10.8 10.1 10.3 10.7 29.0 13.1 9.61 15.8 11.1 9.00 9.11 12.2Nb/U 45.1 39.1 39.3 44.7 40.5 40.5 39.1 41.0 43.7 40.7 40.6 45.5Ba/La 22.0 21.7 21.0 12.4 18.8 16.7 19.6 21.1 25.1 22.7 22.7 23.2Ba/Th 467 433 387 241 333 279 357 380 434 385 389 400(La/Yb)N 78.6 74.2 71.7 75.1 68.5 66.1 69.2 74.9 70.7 65.6 68.0 73.9(Sm/Yb)N 13.8 13.2 13.0 13.4 12.3 12.1 12.3 13.5 12.9 11.8 12.2 13.3

a Reported values for the reference materials are from GeoREM (http://georem.mpch-mainz.gwdg.de/).b Total iron as Fe2O3; Mg# = 100 × Mg2+ / (Mg2+ + Fe2+), assuming Fe3+ / (Fe2+ + Fe3+) = 0.15.

56 Y. Sun et al. / Lithos 208–209 (2014) 53–66

The Re, Ru, Ir, Pt, and Pd remaining in aqua regia were transferred tochloride andfirst separated from thematrix by anion exchange chroma-tography using 2 ml AG 1 × 8, 100–200 mesh resin. Subsequently, Reand Ru were further purified by 0.2 ml AG 1 × 8, 100–200 mesh resin,while Ir–Pt and Pd were further purified by Eichrom®-LN columns.Total procedural blanks were about 2–5 pg for Re, 2 pg for Ir, and

15 pg for Ru, Pt and Pd. The blank corrections were negligible (b1%)for Ir, Ru, Pt and Pd, and were b10% for low-Re samples.

Osmium isotopic compositions were conducted by N-TIMS on a GVIsoprobe-T thermal ionization mass spectrometer at IGGCAS. PurifiedOs was loaded onto platinum filaments with Ba(OH)2 as an ion emitterandmeasured as OsO3

−. The isotopic compositions of Osweremeasured

GLH12-12R GLH12-13 GLH12-14 GLH12-15 GLH12-16 GLH12-17 GLH12-18 GSR-3 -1 Measured GSR-3a Reported GSR-3-2 Measured GSR-3a Reported

48.56 48.08 48.56 48.63 47.83 50.27 50.03 44.72 44.643.03 3.03 3.02 3.00 3.03 3.31 3.24 2.39 2.36

10.77 10.93 10.88 10.87 11.13 11.24 11.22 13.56 13.838.26 8.51 8.58 8.49 8.70 8.20 8.17 13.61 13.400.11 0.11 0.12 0.12 0.12 0.12 0.11 0.17 0.178.65 8.95 9.28 9.10 9.13 7.96 8.17 7.83 7.776.42 6.52 6.46 6.66 6.74 5.54 5.63 8.94 8.812.39 2.64 2.14 2.39 2.57 1.88 2.11 3.27 3.388.54 8.00 8.01 7.93 7.61 9.21 9.17 2.34 2.321.30 1.79 1.34 1.32 1.66 1.09 1.14 0.97 0.951.12 0.62 1.02 1.28 0.76 1.20 0.68 2.48

99.15 99.18 99.41 99.79 99.28 100.02 99.67 100.283.57 3.03 3.74 3.32 2.96 4.90 4.35

70.9 71.0 71.6 71.4 71.0 69.3 70.033.76 31.82 39.83 40.94 34.59 41.38 34.814.34 6.44 6.27 6.48 8.22 3.59 3.65

13.94 12.8 6.59 5.25 8.85 10.87 15.7614.03 16.37 17.32 15.5 18.09 15.22 15.7510.86 10.59 10.88 11.11 11.7 9.15 9.3311.13 9.56 6.05 7.58 6.17 8.38 102.7 5.86 5.03 3.12 5.87 5.78 5.953.08 4.22 3.16 3.11 3.92 2.56 2.68

18.9 18.9 18.9 18.3 18.8 18.5 14.8 15.2338 340 293 299 257 252 140 134210 217 213 192 190 177 141 14043.1 38.0 38.9 36.4 34.2 34.8 49.5 48.6

111 109 112 108 116 114 159 150147 144 144 137 186 166 38.9 37.0

1893 1188 1134 1607 1136 1265 1113 110023.6 22.3 22.8 22.2 23.1 22.7 22.0 22.0

788 759 780 692 843 859 276 27765.9 61.8 64.6 62.6 63.6 63.5 68.8 68.0

3612 3022 3357 3081 5238 4307 532 526152 141 145 137 155 153 58.3 56.0292 271 281 264 289 300 106 10533.3 31.2 32.2 30.5 35.1 34.6 13.3 13.2

113 107 110 104 120 119 54.9 54.018.1 16.8 17.3 16.6 18.7 18.6 10.4 10.24.02 3.72 3.86 3.71 4.10 4.05 3.28 3.20

12.6 11.8 12.2 11.6 12.9 12.5 9.01 8.501.40 1.32 1.36 1.30 1.41 1.38 1.21 1.206.20 5.82 5.99 5.57 6.13 6.05 5.63 5.601.01 0.938 0.966 0.933 0.982 0.973 0.950 0.8802.31 2.15 2.23 2.03 2.24 2.25 2.05 2.000.282 0.269 0.275 0.259 0.268 0.276 0.279 0.2801.51 1.51 1.56 1.41 1.44 1.54 1.50 1.500.208 0.212 0.214 0.196 0.198 0.213 0.183 0.190

20.5 20.1 21.0 18.3 22.7 23.1 6.68 6.503.71 3.46 3.63 3.54 3.58 3.61 4.39 4.30

30.7 28.6 29.9 27.8 30.9 31.3 4.70 4.709.00 8.23 8.66 8.41 8.01 8.26 6.56 6.001.66 1.51 1.50 1.51 0.883 1.56 1.47 1.40

637.8 594.5 614.0 579.5 647.2 653.3

9.51 9.47 9.40 9.51 9.33 9.5639.8 40.9 43.1 41.4 72.0 40.723.8 21.4 23.2 22.5 33.8 28.2

401 367 388 366 654 52172.1 66.9 66.6 70.0 77.0 70.913.3 12.4 12.4 13.2 14.4 13.4

57Y. Sun et al. / Lithos 208–209 (2014) 53–66

either in a peak-jumping mode or a static mode depending on Osconcentrations and sample size. The determined Os isotopic ratioswere corrected for mass fractionation using 192Os/188Os = 3.08271.The in-run precisions for Os isotopic measurements were better than0.2% (2 RSD) for all the samples. The Johnson–Matthey standard of theUniversity of Maryland (UMD) was used as an external standard and

the 187Os/188Os ratio was 0.11378 ± 4 (2SD, n = 20) measured withFaraday cups and 0.1138 ± 4 (2SD, n= 13) with an electron multiplierduring the analytical session in this study.

The measurements of Re, Ir, Pt, Pd and Ru were analyzed at IGGCASon a Thermo-Fisher NEPTUNEMC-ICP-MSwith an electronmultiplier ina peak-jumpingmode or using Faraday cups in a static mode, according

Table 2Sr–Nd–Hf–Pb isotopic compositions of the Xiaogulihe ultrapotassic volcanic rocks.

Sample 87Sr/86Sr 2σm143Nd/144Nd 2σ m εNd(0)a TDM(Nd)b

(Ma)

176Hf/177Hf 2σ m εHf(0)c ΔεHf TDM(Hf)d

(Ma)

206Pb/204Pb 207Pb/204Pb 208Pb/204Pb

GLH12-02 0.705646 0.000008 0.512360 0.000008 −5.4 957 0.282608 0.000003 −5.8 1.4 921 16.468 15.424 36.447GLH12-03 0.705585 0.000008 0.512385 0.000006 −4.9 927 – – – – – 16.443 15.395 36.353GLH12-04 0.705546 0.000010 0.512378 0.000006 −5.1 936 0.282608 0.000003 −5.8 0.9 919 16.470 15.393 36.379GLH12-05 0.705719 0.000010 0.512315 0.000005 −6.3 1008 0.282602 0.000003 −6.0 2.5 928 16.496 15.398 36.411GLH12-06 0.705536 0.000008 0.512366 0.000006 −5.3 956 – – – – 16.496 15.412 36.446GLH12-07 0.705424 0.000010 0.512376 0.000008 −5.1 953 0.282626 0.000003 −5.2 1.6 899 16.514 15.398 36.414GLH12-08 0.705464 0.000011 0.512355 0.000005 −5.5 974 0.282618 0.000003 −5.5 1.9 909 16.520 15.421 36.491GLH12-09 0.705522 0.000008 0.512386 0.000005 −4.9 932 0.282614 0.000002 −5.6 0.8 911 16.474 15.402 36.404GLH12-10 0.705442 0.000010 0.512388 0.000007 −4.9 939 0.282626 0.000003 −5.2 1.2 899 16.535 15.464 36.607GLH12-11 0.705465 0.000008 0.512402 0.000004 −4.6 925 0.282627 0.000003 −5.1 0.8 898 16.552 15.449 36.578GLH12-12 0.705509 0.000008 0.512383 0.000005 −5.0 937 0.282621 0.000003 −5.3 1.2 901 16.479 15.421 36.457GLH12-13 0.705471 0.000007 0.512409 0.000004 −4.5 910 0.282626 0.000003 −5.2 0.5 897 16.503 15.411 36.445GLH12-14 0.705538 0.000008 0.512396 0.000005 −4.7 929 0.282622 0.000003 −5.3 0.8 904 16.500 15.418 36.465GLH12-15 0.705526 0.000008 0.512372 0.000004 −5.2 956 0.282622 0.000004 −5.3 1.5 903 16.494 15.438 36.505GLH12-16 0.705336 0.000010 0.512423 0.000005 −4.2 900 0.282631 0.000003 −5.0 0.3 891 16.496 15.435 36.496GLH12-17 0.705631 0.000008 0.512371 0.000005 −5.2 949 0.282603 0.000003 −6.0 0.9 925 16.486 15.458 36.556GLH12-18 0.705595 0.000008 0.512391 0.000004 −4.8 923 – – – – – 16.460 15.415 36.425BCR-2measured

0.704997 0.000010 0.512648 0.000005 0.282860 0.000004

BCR-2reportede

0.705000 0.512636 0.282878

BHVO-2measured

0.283099 0.000004

BHVO-2reportede

0.283109

a εNd(0) values were calculated using (143Nd/144Nd)CHUR(0) = 0.512638.b TDM(Nd) values were calculated using (147Sm/144Nd)DM(0) = 0.2137 and (143Nd/144Nd)DM(0) = 0.513151.c εHf(0) values were calculated using (176Hf/177Hf)CHUR(0) = 0.282772.d TDM(Hf) values were calculated using (176Lu/177Hf)DM(0) = 0.0384 and (176Hf/177Hf)DM(0) = 0.28325.e Reported values for the reference materials are from GeoREM (http://georem.mpch-mainz.gwdg.de/).

58 Y. Sun et al. / Lithos 208–209 (2014) 53–66

to the measured signal intensity. Mass fractionation was correctedby the Re, Ir, Ru, Pt, and Pd standards that were interspersed amongthe samples. In-run precisions for 185Re/187Re, 191Ir/193Ir, 194Pt/196Pt,105Pd/106Pd, and 99Ru/101Ru were typically 0.1–0.5% (2 RSD). TheUSGS basalt standard BHVO-2 was determined to monitor the accuracyof the analytical procedure, and the obtained values are in goodagreement with the reference values within the analytical errors(see references in Chu et al., 2013) (Table 3).

Table 3Re–Os isotopic compositions of the Xiaogulihe ultrapotassic volcanic rocks.a

Sample Re (ppb) Os (ppb) 187Re/188Os 187Os/188Os

GLH12-04 0.132 0.033 19.6 0.13747GLH12-05 0.096 0.236 1.96 0.12313GLH12-09 0.115 0.137 4.04 0.12558GLH12-10 0.206 0.041 24.2 0.13328GLH12-12 0.170 0.018 45.2 0.14272GLH12-14 0.265 0.065 19.6 0.13429GLH12-15 0.645 0.11874GLH12-16 0.071 0.060 5.73 0.12975DYS12-02 0.046XYS12-04 0.106 0.241 2.13 0.12559JS12-02 0.130 0.037 17.2 0.13814DZS12-03 0.070 0.12488BHVO-2 measured 0.511 0.118 21.1 0.14625BHVO-2 measured 0.544 0.122 21.6 0.14609BHVO-2 reportedc Ref. 1 0.543 0.101

Ref. 2 0.523 0.115

a All the Re, OS, Ir, Ru, Pt, Pd concentrations and the Os isotopic composition data are blank-b γOs values were calculated using (187Os/188Os)CHUR(0) = 0.1270 (Shirey and Walker, 1998c Reported values for the reference material BHVO-2: Ref. 1, Meisel and Moser (2004); Ref.

4. Results

4.1. Whole-rock major and trace elements

Major and trace elemental contents are presented in Table 1. TheXiaogulihe ultrapotassic volcanic rocks have very high K2O contentsranging from 7.61 to 9.29 wt.%, and K2O/Na2O ratios ranging from2.96 to 4.90, mostly higher than 3.0. In terms of SiO2 abundances,

2σm γOsb Ir (ppb) Ru (ppb) Pt (ppb) Pd (ppb) (Pd/Ir)N

0.00038 8.2 0.037 0.057 0.270 0.534 12.00.00014 −3.1 0.173 0.435 0.259 0.299 1.40.00030 −1.1 0.096 0.227 0.260 0.180 1.50.00030 4.9 0.089 0.086 0.322 0.200 1.90.00029 12.4 0.030 0.036 0.262 0.238 6.60.00015 5.7 0.085 0.142 0.207 0.148 1.40.00002 −6.50.00017 2.2 0.075 0.148 0.088 0.122 1.4

0.009 0.015 0.050 0.072 6.40.00023 −1.1 0.255 0.443 0.607 0.432 1.40.00023 8.8 0.061 0.082 0.308 0.229 3.10.00026 −1.70.00005 0.076 0.132 12.0 2.580.00008 0.081 0.138 8.47 2.67

0.058 0.129 10.1 2.940.710 0.123 7.39 2.99

corrected.).2, Shinotsuka and Suzuki (2007).

0

5

10

15K

2O+

Na 2O

(w

t.%)

40 45 50 55 60

Tephri-phonolite

FoiditePhono-Tephrite

TephriteBasanite

Trachy-basalt

Basaltictrachy-andesite

Trachy-andesite

Basalt

Alkaline

Subalkaline

KeluoGLH

EKSWDLC WEK

SiO2 (wt.%)

Fig. 3. Total alkalis vs. SiO2 for the Xiaogulihe ultrapotassic volcanic rocks. Rock classifica-tion is after Le Bas et al. (1986). Xiaogulihe (GLH) data from this study; Keluo data fromZhang (1992) and Zhang et al. (1995); WDLC and EKS data from Chu et al. (2013).

10 12 14 168

10

12

14

16

18

20

NE China

2 4 6 8

2 4 6 8 10 12 14 16

CaO

(w

t.%)

2

4

6

8

10

12

14

a

b

Fe/

Mn

50

60

70

80

90

100

Kelu o xenoliths

Hawaii island

c

KeluoGLH

EKSWDLC WEK

Al 2O

3 (w

t.%)

MgO (wt.%)

MgO (wt.%)

59Y. Sun et al. / Lithos 208–209 (2014) 53–66

these ultrapotassic rocks display a narrow range from 47.8 to 50.3 wt.%(Table 1), while theWEK potassium-rich volcanic rocks range from 42.3to 53.5 wt.% (Chu et al., 2013; Shao et al., 2009). In addition, theseultrapotassic rocks fall in the tephri-phonolite and phono-tephritefieldsin a TAS plot (Fig. 3) and in the ultrapotassic series in a K2O vs. Na2Odiagram of Foley et al. (1987) (Fig. 4). Meanwhile, these ultrapotassicrocks have relatively low Al2O3 and CaO with a given MgO contentcompared to the other Cenozoic basalts in NE China (Fig. 5a and b).The Mg# values of these ultrapotassic rocks range from 69.1 to 71.6,falling within those of primary magma (68–75, Frey et al., 1978).

The Xiaogulihe ultrapotassic rocks are characterized by a noticeableenrichment in light rare earth elements (LREEs) relative to heavyrare earth elements (HREEs) ((La/Yb)N = 65.6–78.6) in a chondrite-normalized REE diagram (Fig. 6a; Table 1). The HREEs are also stronglyfractionated with (Sm/Yb)N = 11.8–14.4 (Table 1). The similarprimitive mantle-normalized trace element patterns indicate that theserocks are significantly enriched in LILEs with obviously positiveanomalies of K and Pb, and depleted in HFSEs with apparently Th, U,Nb, Ta, and Ti negative anomalies (Fig. 6b). In addition, the Ce/Pb ratios(9.0–15.8) of these ultrapotassic rocks are lower than those of averageMORB and OIB (25 ± 5, Hofmann et al., 1986), while the Nb/U ratios(39.1–45.5) are similar to those of average MORB and OIB (47 ± 10,Hofmann et al., 1986) (see Fig. 7a and b).

UltrapotassicRocks

Potassic Rocks

K2O/Na2O=3

K2O/Na2O=2

K2O/Na2O=1

0 1 2 3 4 5 6 70

2

4

6

8

10

12

K2O

(w

t.%)

KeluoGLH

EKSWDLC WEK

Na2O (wt.%)

Fig. 4. K2O vs. Na2O plot for the Xiaogulihe ultrapotassic volcanic rocks. The sources of thedata are the same as those for Fig. 3.

Mg #75 80 85 90 95 100

30

40

Fig. 5. Whole-rock CaO vs. MgO (a) and Al2O3 vs. MgO (b), and Fe/Mn vs. Mg# ofphenocryst olivine (c) for the Xiaogulihe ultrapotassic volcanic rocks. In (a) and (b):GLH data are from this study; data for the Keluo potassic volcanic rocks are from Zhang(1992) and Zhang et al. (1995); data for the WDLC and EKS potassic volcanic rocks andother Cenozoic NE China basalts are from Basu et al. (1991), Liu et al. (1994), Zhanget al. (1995), Zou et al. (2003), Chen et al. (2007), Yan and Zhao (2008) and Chu et al.(2013). In (c): phenocryst olivine data for GLH are from this study; the Keluo mantlexenolith data are from Zhang et al. (2011); field for the Hawaii island basalts is based ondatabase GEROC (http://georoc.mpch-mainz.gwdg.de/georoc/).

4.2. Whole-rock Sr–Nd–Hf–Pb isotopes

Strontium, Nd, Hf and Pb isotopic compositions of the Xiaoguliheultrapotassic rocks are presented in Table 2, and illustrated in Figs. 8

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Roc

k/C

hond

rite

1

10

100

1000

OIB

N-MORB

GLH

Keluo

EKS

WDLC

a

RbBa

ThU

NbTa

KLa

CePb

PrSr

PNd

ZrHf

SmEu

TiGd

TbDy

YHo

ErTm

YbLu

Roc

k/P

rimiti

ve M

antle

0.1

1

10

100

1000

10000

OIB

N-MORB

Keluo

EKS

WDLC

b

Fig. 6. Chondrite-normalized REE (a) and primitive mantle-normalized trace elementpatterns (b) for the Xiaogulihe ultrapotassic volcanic rocks. Chondrite and primitiveman-tle data are after Sun andMcDonough (1989). Field for the Keluo potassic volcanic rocks isbased on Zhang et al. (1995); fields forWDLC and EKS are from Chu et al. (2013); data forOIB and N-MORB are from Sun and McDonough (1989).

0 30 180 210 240 270 300 330 360

Ce/

Pb

0

5

10

15

20

25

30

35

MORB & OIB

LCC

UCC a

0 20 40 60 80 100

Nb/

U

0

20

40

60

80

LCC

UCC

MORB & OIB

b

Ba/Th0 100 200 300 400 500 600 700

Ba/

La

0

10

20

30

40

50

60

GaussbergLeucite Hills

K-hollandite

c

KeluoGLH

EKSWDLC WEK

Ce(ppm)

Nb(ppm)

60 90 120 150

Fig. 7. Ce/Pb vs. Ce (a), Nb/U vs. Nb (b) and Ba/La vs. Ba/Th (c) diagrams for the Xiaoguliheultrapotassic volcanic rocks. In (a) and (b): data for the Keluo potassic rocks are fromZhang (1992) and Zhang et al. (1995); data for the WDLC and EKS potassic rocks arefrom Chu et al. (2013); data for the upper (UCC) and lower (LCC) continental crusts arecalculated after Rudnick and Gao (2003); and data for MORB and OIB are from Hofmannet al. (1986). In (c): the data sources for Keluo, WDLC and EKS are the same as (a) and(b); data for other ultrapotassic igneous rocks are fromMurphy et al. (2002) andMirnejadand Bell (2006); data for K-hollandite are from Rapp et al. (2008).

60 Y. Sun et al. / Lithos 208–209 (2014) 53–66

and 9 along with the WEK potassic volcanic rocks and worldwideultrapotassic igneous rocks for comparison.

The Xiaogulihe ultrapotassic rocks show restricted variations in bothSr and Nd isotopic ratios with 87Sr/86Sr varying from 0.7053 to 0.7057,and εNd from−4.2 to−6.3 (Fig. 8a and b; Table 2). These ultrapotassicrocks alongwith theWEK potassic rocks show a ‘steep trend’ (relativelylow and restricted 87Sr/86Sr ratios, and low εNd values) defined by thewell known NW American ultrapotassic igneous rocks (e.g. SmokyButte, Leucite Hills and Crazy Mountains) and Aldan Shield lamproites,whereas Mediterranean lamproites (e.g. Italy, Spain, Serbia,Macedonia) and Western Australia lamproites form another evolvingtrend (relatively high and unrestricted 87Sr/86Sr ratios, and low εNdvalues). The TDM (Nd) model ages of these ultrapotassic rocks rangefrom 900 to 1008 Ma, which are regarded as the minimum ages ofsource metasomatism.

The Xiaogulihe ultrapotassic rocks have very homogeneousHf isoto-pic compositions, with 176Hf/177Hf ratios ranging from 0.28260 to0.28263 (Fig. 8c and d; Table 2). Similar to the Nd model ages, theseultrapotassic rocks give uniform TDM (Hf) model ages varying from891 to 928 Ma. In order to evaluate the degree of Nd–Hf isotopicdecoupling, we have calculatedΔεHf values, whereΔεHf is the differencein εHf relative to the εNd–εHf terrestrial array, defined as ΔεHf = εHf −1.55εNd − 1.21 (Vervoort et al., 2011). The ΔεHf values of theseultrapotassic rocks range from 0.3 to 2.5, indicating no obvious Nd–Hfisotopic decoupling (Table 2).

The Xiaogulihe ultrapotassic rocks show extremely unradiogenic Pbisotopic compositions with 206Pb/204Pb ranging from 16.44 to 16.55,207Pb/204Pb from 15.39 to 15.46, and 208Pb/204Pb from 36.35 to 36.61.They plot on the left side of the ‘Geochron’ and upon the NorthernHemisphere Reference Line (NHRL, Hart, 1984), and show a trendtowards the LOMU2 end-member defined by Chen et al. (2007) (Fig. 9).

0.705 0.710 0.715 0.720 0.725-30

-20

-10

0

10

20

Leucite Hills

Western Australia

Southern Afican

Western Australia

Crazy Mt.

Smoky Butte

Mediterranean

Aldan Shield

Gaussberg

DMM

FOZO

HIMU

LOMU 1

LOMU 2

EMI

EMII

a

KeluoGLH

EKSWDLC WEK

ε Nd(

t)

εNd(t)

εNd(t)

0.7045 0.7050 0.7055 0.7060 0.7065-10

-8

-6

-4

-2

0

2

LOMU 2

EMI

b

ε Nd(

t)

-6 -5 -4 -3 -2-12

-10

-8

-6

-4

-2

0

2

Terrestrial Array

d

-20 0 5-40

-30

-20

-10

0

10

20

30

MORB

OIB

Marinesediments

Transitionalkimberlites

Group I kimberlites

Aldan Shield

Macedonia

Serbia

Group IIkimberlites

ItalySpain

Leucite Hills

Terrestrial Array c

ε Nf(t

)ε N

f(t)

87Sr/86Sri

87Sr/86Sri

-15 -10 -5 10 15

Fig. 8. Sr–Nd (a) and (b), and Nd–Hf (b) and (c) isotope correlation diagrams for the Xiaogulihe ultrapotassic volcanic rocks. In (a) and (b): Keluo data are from Zhang et al. (1995);WDLCand EKS data fromBasu et al. (1991), Zhang et al. (1995), Fan et al. (2001), Zou et al. (2003) and Chu et al. (2013); LOMU1, LOMU2, DMM, FOZO, HIMU, EMI, and EMII endmembers can befound in Chen et al. (2007); data for other ultrapotassic igneous rocks are fromMcCulloch et al. (1983), Vollmer et al. (1984), Fraser et al. (1985), Dudas et al. (1987), O'Brien et al. (1995),Murphy et al. (2002), Prelevic et al. (2004, 2005, 2008) andDavies et al. (2006). In (c) and (d):WDLC and EKS data are fromChu et al. (2013);MORB andOIB fields are based on theGEROCdatabase (http://georoc.mpch-mainz.gwdg.de/georoc/); field for marine sediments is based on Vervoort et al. (2011); data for kimberlites and other ultrapotassic igneous rocks are fromSalters and Hart (1991), Nowell et al. (2004), Davies et al. (2006) and Prelevic et al. (2010); the terrestrial array is based on Vervoort et al. (2011).

61Y. Sun et al. / Lithos 208–209 (2014) 53–66

4.3. Re–Os isotopes and PGE

The Re–Os isotopic compositions of the Xiaogulihe ultrapotassicrocks are given in Table 3. In general, the Xiaogulihe ultrapotassicrocks have higher Re (0.096–0.265 ppb) and lower Os (0.018–0.236 ppb) abundances compared with the nearby WDLC and EKSpotassic rocks (Re: 0.079–0.130 ppb; Os: 0.067–0.389, Chu et al.,2013). The 187Os/188Os ratios of these ultrapotassic rocks, varying from0.1187 to 0.1427, are generally lower than those of the WDLC and EKSpotassic rocks (0.13–0.17, Chu et al., 2013), and the continental basaltsfrom Hannuoba (Jiang and Zhi, 2010) and Emeishan (EMS) (Xu et al.,2007) (Fig. 10a). The sample GLH12-15, with the lowest 187Os/188Osratio (0.1187) and the highest Os content (0.645 ppb) among all ofthe analyzed samples, is considered to have contained a small amountof mantle materials such as xenocrysts that incorporated primarysulfides or PGE micro-alloys during magma ascent through the over-lying ancient SCLM (Chu et al., 2013). In addition, these ultrapotassicrocks have broadly the same 187Os/188Os ratios with the mantlexenoliths hosted in the nearby Keluo potassic rocks (0.1146–0.1319,Zhang et al., 2011). These relatively low 187Os/188Os ratios possiblysuggest a lithospheric origin.

The PGE abundances of these ultrapotassic rocks are presented inTable 3. The Xiaogulihe ultrapotassic rocks have low PGE abundances,

ranging from 0.030 to 0.173 ppb for Ir, 0.036 to 0.435 ppb for Ru,0.088 to 0.322 ppb for Pt, and 0.122 to 0.534 ppb for Pd, similar tothose of the nearby WDLC and EKS potassic rocks (Chu et al., 2013).As shown in Fig. 10b, these ultrapotassic rocks display slight PPGEenrichment compared to IPGE ((Pd/Ir)N = 1.4–1.9, Table 3), which isconsidered as the feature of lithospheric mantle.

5. Discussion

Asmentioned above, themain problems concerning the petrogenesisof the Xiaogulihe ultrapotassic volcanic rocks can be summed up intothree issues: (1) the site of ultimate mantle source; (2) the source andgeochemical characteristics of the metasomatic fluids and/or melts;and (3) the dominant source rocks of these ultrapotassic rocks.

5.1. Crustal contamination

Before these ultrapotassic volcanic rocks can be used to constraintheir ultimate mantle source, the influence of shallow processes suchas crustal contamination should be evaluated. Significant contaminationof the WEK potassium-rich rocks by the upper continental crusthas been revealed by previous studies based on trace elements andSr–Nd–Pb isotopic compositions (Zhang, 1992; Zhang et al., 1998).

16.0 17.0 18.0 19.0

207 P

b/20

4 Pb

208 P

b/20

4 Pb

206Pb/204Pb

206Pb/204Pb

EMI

LOMU 2

LOMU 1

DMM

EMI I

Western Australia Mediterranean

Aldan Shield

Smoky Butte

Crazy Mt.

Leucite Hills

Gaussberg

Southern Afican

FOZO

Geo

chro

n

NHRL

Recent sediments

0.5 Ga

1.0 Ga

1.5 Ga

2.0 Ga

a

16.0 17.0 18.0 19.0

Recent sediments

0.5 Ga

1.0 Ga

1.5 Ga

2.0 Ga

LOMU 1

LOMU 2

EMI EMII

DMM

FOZO

Leucite Hills

Smoky Butte

Western Australia

Southern AficanCrazy Mt.

Mediterranean

Gaussberg

Aldan Shield

NHRL

b

KeluoGLH

EKSWDLC WEK

15.8

15.6

15.4

15.2

39.5

38.5

37.5

36.5

35.5

Fig. 9. Pb isotope correlation diagrams for theXiaogulihe ultrapotassic volcanic rocks. Datafor Keluo are from Zhang (1992), Zhang et al. (1995) and Shao et al. (2008); WDLC andEKS data are from Basu et al. (1991), Zhang (1992), Zhang et al. (1995) and Zou et al.(2003); the data sources for mantle end members are the same as Fig. 8(a) and(b); data for other ultrapotassic igneous rocks are from Fraser et al. (1985), Dudas et al.(1987), O'Brien et al. (1995), Murphy et al. (2002), Prelevic et al. (2004, 2005, 2008)andDavies et al. (2006). The Pb isotope evolutionmodel for continental crust-derived sed-iments is based on Kuritani et al. (2011). NHRL is after Hart (1984).

187 O

s/18

8 Os

187 O

s/18

8 Os

Crustal

contamination

Fractionalcrystallization

HannuobaEMS

a

Roc

k/C

hond

rite

GLHPrimitive mantle

MORB

OIB

UCC

b

0.11

0.12

0.13

0.14

0.16

0.18

0.20

R2=0.86

R2=0.47

3.5%

1%

Keluo xenoliths

c

0 60

0

0.15

0.17

0.19

GLH

EKSWDLC

WEK

1/Os

Os(ppb)

0.7

0.6

0.5

0.4

0.3

0.2

0.10.80.70.60.50.40.30.20.1

10-1

10-2

10-3

10-4

10-5

10-6

Os Ir Ru Pt Pd Re

10 20 30 40 50

Fig. 10. 187Os/188Os vs. Os concentration (a), chondrite-normalized PGE patterns (b) and187Os/188Os vs. 1/Os (c) diagrams for the Xiaogulihe ultrapotassic volcanic rocks. In(a): data for WDLC and EKS are from Chu et al. (2013); Hannuoba data are from Jiangand Zhi (2010); and EMS data are from Xu et al. (2007). In (b): chondrite values arefrom McDonough and Sun (1995); primitive mantle (PM) values are from Becker et al.(2006); OIB and MORB values are from Day (2013) and references therein; and UCC(upper continental crust) values are from Rudnick and Gao (2003). In (c): WDLC andEKS data sources are the same as Fig. 10a; the Keluo mantle xenolith data are fromZhang et al. (2011). The dashed lines represent binary mixing lines (see details in text).

62 Y. Sun et al. / Lithos 208–209 (2014) 53–66

However, as these ultrapotassic rocks are strongly enriched in lithophileelements, the Sr–Nd–Pb isotopic systems are not effective to identify asmall amount of lower continental crust addition. In this regard, theRe–Os isotopic system can easily detect the minor effect of a lessevolved lower continental crust which generates 187Os/188Os ratiosthat are distinct from those of the mantle source.

Previous studies have suggested that the samples with low Os con-tents and abnormally high 187Os/188Os ratios most likely experiencedcontamination of crustal materials which have radiogenic Os isotopiccompositions (Reisberg et al., 1993; Widom, 1997). In the 187Os/188Osvs. Os concentration variation diagram (Fig. 10a), continental floodbasalts such as Hannuoba (Jiang and Zhi, 2010) and EMS (Xu et al.,2007) display a relatively complex initial 187Os/188Os ratios with Oscontents. Those with low Os contents but anomalously high initial187Os/188Os ratios were considered to have experienced shallow-levelcrustal contamination.

The Xiaogulihe ultrapotassic rocks show a relatively narrow range in187Os/188Os ratios despite awide range inOs concentrations as shown inFig. 10a and Table 3. It ismost likely that these ultrapotassic samples hadnot been significantly affected by crustal materials during magmaascent. We assume an Os isotopic ratio of ca. 0.128, but varying Os con-tents (0.018–0.645 ppb, depended on samples analyzed here, Table 3)as the primitive ultrapotassic magma of these rocks when they arrivedat the lower continental crust. The extent of crustal contamination of

these ultrapotassic rocks can be evaluated through a simple binarymixing model when an Os isotopic ratio of 0.8 and Os content of0.049 ppb are used to represent the lower continental crust (Chesleyet al., 2002; Saal et al., 1998). It is found that the addition of less than1% of the lower continental crust can approximately replicate theobserved correlated variation between 187Os/188Os and 1/Os for the

63Y. Sun et al. / Lithos 208–209 (2014) 53–66

Xiaogulihe samples, whereas that of the WDLC samples needs theaddition of about 3.5% of the lower crustal materials (Fig. 10c). Afterall, addition of such a minor amount of lower crustal materials wouldnot have any effect on incompatible elements and Sr–Nd–Hf–Pb isotopicratios for the LILE- and LREE-rich Xiaogulihe ultrapotassic rocks.

The high Mg# values (69.1–71.6, Table 1) of these ultrapotassicrocks and the presence of mantle xenoliths included in the nearbyKeluo, WDLC and EKS potassic rocks (Zhang et al., 2000, 2011) alsosuggest a rapid eruption and the effect of crustal contamination isnegligible. Thus, we confirm that the geochemical characteristics ofthe Xiaogulihe ultrapotassic rocks can be reliably used to trace theirmantle source.

5.2. Ultimate mantle source

Due to the complex and varying structure of themantle beneath thisregion, several competing theories have been put forward concerningthe ultimate mantle source of the WEK potassium-rich rocks. Based onthe high enrichment in K2O and lithophile elements (LREEs and LILEs),the strong fractionated REEs, significant 230Th excesses, and EMI-likeisotopic characteristics, Zhang et al. (1995), Zou et al. (2003) and Chuet al. (2013) suggested that the WEK potassium-rich rocks directlyoriginated from the SCLM by low-degree partial melting (5–7%) ofphlogopite-bearing garnet peridotite at depths of 80–120 km. Inaddition, based on the high SiO2 contents relative to incompatibletrace elemental concentrations of the WEK potassium-rich volcanicrocks, low 3He/4He ratios of the ultramafic xenoliths enclosed in thesepotassium-rich rocks, and the seismic tomographic images, Chen et al.(2007) proposed that these potassium-rich rocks were most likelyproduced by low-degree partial melting of shallow SCLM (absence ofgarnet) in the presence of phlogopite and/or amphibole.

On the other hand, several studies have suggested that late Cenozoicbasalts throughout eastern Asia mainly originated from shallowasthenosphere characterized by two distinct, broad domains—a DMM-EMI mixing array for NE China, and a DMM-EMII mixing array for SEAsia (Choi et al., 2006; Xu et al., 2005; Zou et al., 2000). Furthermore,on the basis of the high temperature of the magma shortly beforeeruption, Kuritani et al. (2013) argued that both the potassic- andEMI-like signatures of the WEK potassic volcanic rocks originatedfrom the mantle transition zone. The transition zone source had beenmetasomatized by potassium-rich sediment-derived fluids ~1.5 Gaago through the stagnation of an ancient slab (Kuritani et al., 2011;Murphy et al., 2002), and the recent sediments and peridotites in thestagnant Pacific slab have also played an important role in the sourcemodification.

0.7048 0.7050 0.7052 0.7054 0.7056 0.7058

206 P

b/20

4 Pb

KeluoGLH

WDLC WEK

87Sr/86Sr

17.2

17.0

16.8

16.6

16.4

Fig. 11. 206Pb/204Pb vs. 87Sr/86Sr for the Xiaogulihe ultrapotassic volcanic rocks. Data forthe Keluo potassic volcanic rocks are from Zhang et al. (1995); WDLC data are after Basuet al. (1991), Zhang et al. (1995) and Zou et al. (2003).

According to this study, the Xiaogulihe ultrapotassic rocks showextreme enrichments in potassium and lithophile elements (LREEsand LILEs) (Fig. 6), particularly highly fractionated REE (Fig. 6a), andEMI-like Sr–Nd–Hf–Pb isotopic characteristics (Figs. 8 and 9). All thesegeochemical characteristics are similar to those of previous studies forthe WEK potassic rocks (Chu et al., 2013; Zhang et al., 1995; Zou et al.,2003), indicating a SCLM source for the potassium-rich magmas.Furthermore, the Xiaogulihe ultrapotassic rocks have relatively low187Os/188Os ratios and flat chondrite-normalized PGE patterns(Fig. 10), which are considered as the characteristics of the SCLM. TheSCLM typically has lower Os isotopic compositions than the primitiveupper mantle (e.g. 187Os/188Os ratio of 0.1296, Meisel et al., 2001), andthe plume-related OIB-type mantle (e.g. 187Os/188Os ratios rangingfrom 0.13 to 0.15, Shirey and Walker, 1998). Meanwhile, the negativecorrelation between 87Sr/86Sr and 206Pb/204Pb (Fig. 11) in most of thepotassium-rich rocks with EMI signatures is also attributed to thepresence of a potassic phase, mostly phlogopite, in the SCLM resultingfrom mantle metasomatism, a deep process that is restricted only tothe lithosphere. The presence of phlogopite in some Keluo and WDLClherzolite xenoliths (Zhang et al., 2000, 2011) also argues that theSCLM beneath the WEK volcanic field had been metasomatizedby potassium-rich melts. Moreover, significant 230Th excesses in thenearbyWDLC and EKS potassic rocks (Zou et al., 2003), strong fraction-ation of the REEs, and no obvious Nd–Hf isotopic decoupling for all theWEK potassium-rich rocks indicate that garnet was mainly retained inthe mantle source. In other words, the melting degree of their mantlesource is relatively low that the geochemical features of garnet in theSCLM have not been exported into the resulting potassium-richmagma. Therefore, it is most likely that the Xiaogulihe ultrapotassicrocks originated from phlogopite-bearing garnet source rocks withinthe SCLM.

Alternatively, Choi et al. (2006) proposed that the WEK potassium-rich volcanic rocks mainly originated from shallow asthenospherewhich hosted ancient SCLM that delaminated from eastern China.Let alone whether the phlogopite and/or amphibole serving as potassi-um holders are stable or not under such high P–T conditions, the Redepletion ages (TRD ~ 1.9–2.0 Ga) of the mantle xenoliths entrained inthe potassium-rich volcanic rocks argue against the delamination ofthe SCLM in this area (Zhang et al., 2011). On the other hand, Choiet al. (2008) and Chu et al. (2009) have argued that the delaminatedArchean SCLM beneath eastern China does not have EMI-likecompositions. Therefore, it is difficult to explain the potassium-richvolcanism in NE China by the delamination model.

Recently, Kuritani et al. (2013) proposed that the potassium-richvolcanic rocks mainly originated from the mantle transition zonewhich had beenmodified by an ancient slab and then by a recent Pacificslab.However, several studies have argued that the recent subduction ofthe western Pacific plate might not directly contribute subduction-related fluids to the source rocks of these potassium-rich volcanicrocks (Zhang et al., 1995; Zou et al., 2003). Because of the great distance(N2000 km) between the WEK potassium-rich volcanic field and theJapan Arc, the subducted slab may have lost fluids released fromsubducted sediments before it reached the study area. Moreover, themodification of source rocks by sediment-derived fluids in the last350 ka was also ruled out due to the apparent 230Th excesses in theWDLC potassic rocks (Zou et al., 2003). The basic argument whichsupported the mantle transition zone model was that the estimatedtemperature (1250 °C) of these potassium-rich magmas was higherthan the maximum temperature (1180 °C) of the lithospheric mantlebeneath theWEK volcanic field (Kuritani et al., 2013). In the thermody-namicmodel of Kuritani et al. (2013), the temperature shows a negativecorrelation with equilibrium Mg# of olivine phenocrysts for thepotassium-rich magma. The Mg# values (85.2 and 86.1) of olivinephenocrysts of relatively evolved WDLC potassic rocks were used tocalculate the magma temperature. However, our study shows that theolivine phenocrysts of the Xiaogulihe samples haveMg# values ranging

64 Y. Sun et al. / Lithos 208–209 (2014) 53–66

from86.3 to 89.5, generally higher than those of theWDLC samples. Thecalculated temperature is quite lower than that of Kuritani et al. (2013),and falls within the temperature ranges of the SCLM beneath the WEKpotassium-rich volcanic field.

Consequently, it is most likely that the primaryWEK potassium-richmagma was generated by a low-degree partial melting of phlogopite-bearing garnet source rocks within the lower SCLM at depths of80–120 km which had been metasomatized by potassium-rich silicatemelts (Chu et al., 2013; Zhang et al., 1995, 2000; Zou et al., 2003).

5.3. Source of the metasomatic potassium-rich silicate melts

As argued above, the Xiaogulihe ultrapotassic rocks mainlyoriginated from the lower SCLM that had been metasomatized bypotassium-rich silicate melts. Therefore, the unusually geochemicalcharacteristics of these ultrapotassic rocks, such as high K2O contents,abnormally unradiogenic Pb isotopic compositions, and relatively low87Sr/86Sr ratios can beused to trace the source of themetasomaticmelts.

5.3.1. Extremely high K2O contentsRecently, Chu et al. (2013) argued that the potassic characteristic

of the WEK potassium-rich volcanic rocks could not be owed tometasomatism by melts generated from low-degree partial melting ofdeep asthenosphere (Zhang et al., 2000) based on the isotopic com-positions, as all potassium-rich rocks fall outside the range of the OIBand MORB. They instead suggested that the original ancient lowercontinental crust together with SCLM beneath the WEK volcanic fieldhad delaminated into the asthenosphere during lithospheric thinningin eastern China. The newly accreted SCLM had been metasomatizedby potassium-rich silicate melts derived from the delaminated lowercontinental crust and then generated the WEK potassium-rich magma.However, we should realize that the average K2O content of the lowercontinental crust is only about 0.61 wt.% (Rudnick and Gao, 2003). It isdifficult to imagine that the lower continental crust with such a lowK2O content could modify the newly accreted SCLM and then producethe potassium-rich magma with extremely high K2O contents (7.61 to9.29 wt.%, Table 1). Furthermore, as mentioned above, lithosphericdelamination did not occur in this area (Zhang et al., 2011). Therefore,it is unlikely that the ultimate mantle source of the Xiaoguliheultrapotassic rocks had been metasomatized by potassium-rich silicatemelts derived from the delaminated lower continental crust.

On the other hand, previous studies have suggested that thepotassium-rich silicate melts might possibly originate from thesubducted continental-derived sediments that have recycled into andaccumulated in the mantle transition zone (Kuritani et al., 2011, 2013;Murphy et al., 2002; Rapp et al., 2008). Mantle tomographic studiesalso show that oceanic lithospheric slabs containing a considerableamount of subducted sediments can accumulate in the mantletransition zone (Simons et al., 1999). Under mantle transition zoneconditions, the followingminerals derived from continental sedimenta-ry protoliths have been observed: majorite, K-hollandite, stishovite,calcium aluminum silicate, and Ca-perovskite (Irifune et al., 1994). K-hollandite in subducted sediments is considered as the only majorphase which can release a significant amount of K and incompatibleelements such as Ba and Pb through a breakdown under certainconditions (Rapp et al., 2008).

The positive spikes of Ba, K, and Pb are observed in the trace elementconcentration patterns of the Xiaogulihe ultrapotassic rocks along withother potassium-rich rocks (Fig. 6b). The composition of K-hollandite(Rapp et al., 2008) seems to lie on the extension of the compositionalvariation of the Gaussberg and Leucite Hills ultrapotassic rocks in theBa/La–Ba/Th diagram (Fig. 7c), indicating the contribution of a sedimentcomponent in their magma sources. Nevertheless, the other potassium-rich volcanic rocks fall outside the trend. As suggested by Plank andLangmuir (1998), the isotopic and chemical compositions of sedimentsvary inmodern subduction zones. By analogy, the isotopic and chemical

compositions of sediments with different ages are also possibly thoughtto have varied. Moreover, K-hollandite primarily retains the incom-patible element characteristics of its host rock. Thus, the compositionof K-hollandite (Rapp et al., 2008) could only represent specificsubducted sediments that have altered magma sources of theGaussberg and Leucite Hills ultrapotassic rocks. The different K2Ocontents of the worldwide potassium-rich igneous rocks can alsobe explained by the various contributions of a specific sedimentcomponent in the mantle sources, as suggested by the diverse Ba/Thand Ba/La ratios (Fig. 7c).

Therefore, it is most likely that the potassium-rich silicate meltswere generated by partial melting of subducted continental-derivedsediments that contained K-hollandite as the major K and incompatibleelements such as a Ba and Pb reservoir.

5.3.2. Extremely unradiogenic Pb isotopic compositionsThe Xiaogulihe ultrapotassic rocks have extremely unradiogenic Pb

isotopic compositions (206Pb/204Pb = 16.44–16.55, Fig. 9 and Table 2),which require the isolation of a low-μ mantle source with low initialPb isotopic ratios for periods of greater than 1.5 Ga. Therefore, themetasomatic potassium-rich silicate melts may have the similarlylow-μ signature and low initial Pb isotopic ratios.

It is well known that the partial melting of peridotite cannot cause alarge U/Pb fractionation. Hence, the melts produced by low-degreepartial melting of deep asthenosphere cannot have a significantly low-μ signature. Even if the melts derived from deep asthenosphere havelow initial Pb isotopic ratios, the long-term residencewould significant-ly alter the Pb isotopic compositions. On the other hand, the youngsubducted sediments generally have radiogenic Pb isotopic composi-tions (206Pb/204Pb = ~18.8, recent Pacific sediments, Kuritani et al.,2011). The extremely unradiogenic Pb isotopic compositions of theXiaogulihe ultrapotassic rocks indicate that at least the Pb isotopicsignatures of the ultimate mantle source of these ultrapotassic rockscannot be attributed to the modifications by Phanerozoic subductionevents (Zhang et al., 2000).

Alternatively, slab dehydration driven by metamorphism at thetime of subduction is a plausible mechanism for fractionation ofparent/daughter ratios (Becker et al., 2000). If the high oxygen fugacityprevailed during subduction; oxidation of U4+ to the more soluble U6+

could have caused U to be lost relative to Pb, thereby resulting in low-μ(Murphy et al., 2002). Moreover, the ancient subducted sediments haverelatively lower Pb isotopic compositions than their modern counter-parts (206Pb/204Pb ~ 16.0, 2.0 Ga; 206Pb/204Pb ~ 16.8, 1.5 Ga, Fig. 9).Therefore, the ancient subducted sediments (N1.5 Ga) accumulated inthe mantle transition zone would be a plausible source of the meta-somatic potassium-rich silicate melts for the Xiaogulihe ultrapotassicrocks. As shown in Fig. 9a, the worldwide potassium-rich igneousrocks mostly plot to the left of the ‘Geochron’, showing a large rangein both 207Pb/204Pb and 206Pb/204Pb isotopic ratios. These isotopicvariations can be explained if the ages of the subducted continental-derived sediments are varied and if small allowances are made for thewide range in the 207Pb/206Pb ratios of the upper continental crust(Murphy et al., 2002). The rare ultrapotassic rocks that plot to theright of the ‘Geochron’ are from the Mediterranean (Prelevic et al.,2008). They are explained as the ultimate lithospheric mantle sourcewhich has been altered by melting products of continental-derivedsediments that were very recently subducted into the mantle (Fig. 9).This seems to be consistent with their occurrence close to the Alpinefront, along which sediments were subducted throughout the recentclosure of the Tethys ocean.

5.3.3. Relatively low 87Sr/86Sr ratiosThe Sr and Nd isotopic characteristics of the global ultrapotassic

igneous rocks are also highly variable and show two distinct trends inthe 87Sr/86Sr vs. εNd diagram (Fig. 8a). Ultrapotassic rocks from theMediterranean and Western Australia plot on a trend of increasing

65Y. Sun et al. / Lithos 208–209 (2014) 53–66

87Sr/86Sr, whereas ultrapotassic rocks from Xiaogulihe, Smoky Butte,Leucite Hills, Aldan Shield, and Crazy Mountains plot on the othertrend of a decreasing εNd with relatively narrow change in 87Sr/86Srratios. As demonstrated above, the potassium-rich silicate melts maybe generated by partial melting of subducted continental-derivedsediments. Therefore, the variations possibly represent heterogeneitiesin the continental-derived sediments being subducted into the mantleand more importantly in subduction zone processes (Murphy et al.,2002). It has been suggested that the sources for the Mediterraneanand Western Australian ultrapotassic rocks have evolved with highRb/Sr ratios resulting from the dominant K-bearing phases over Ca-bearing phases in the continental-derived sediments during theirsubduction into the deep mantle. We propose here that the mantlesource domains for the Xiaogulihe, Smoky Butte, Leucite Hills, AldanShield, and Crazy Mountains ultrapotassic rocks probably have evolvedwith low Rb/Sr ratios resulting from the metasomatized melts derivedfrom the ancient subducted sediments.

5.4. Source lithology: peridotite vs. pyroxenite

As argued above, the Xiaogulihe ultrapotassic volcanic rocks mainlyoriginated from phlogopite-bearing garnet source rocks within thelower SCLM. However, phlogopite-bearing garnet source rocks couldbe garnet pyroxenites or garnet peridotites.

If garnet pyroxenites are the dominant source rocks, the partialmelts would be enriched in Al2O3 and particularly CaO (see referencesin Zou et al., 2003). Nevertheless, the Al2O3 and CaO abundancesfor the Xiaogulihe ultrapotassic rocks are relatively lower than otherCenozoic basalts in NE China at a given MgO (Fig. 5). Moreover,Sobolev et al. (2005, 2007) has argued that pyroxenite-derived meltswould have higher Fe/Mn ratios than those of peridotite-derivedmelts. The involvement of olivine-free pyroxenites can be recognizedfrom the presence of olivine phenocrysts with high Fe/Mn ratios.However, as shown in Fig. 5c and Table S1, olivine phenocrysts ofthe Xiaogulihe ultrapotassic rocks mostly have Fe/Mn ratios b 60,even lower than those crystallized from the peridotite-derived melts(~60–70) as estimated by Herzberg (2011). Kelley and Cottrell (2009)has suggested that those olivines with Fe/Mn ratios b 60 might crystal-lize frommore oxidizedmelts. Therefore, we propose that the dominantsource rocks for the Xiaoguliheultrapotassic rocks are garnet peridotitesrather than garnet pyroxenites.

6. Concluding remarks

The following conclusions can be reached based on the elementaland isotopic studies on the Xiaogulihe ultrapotassic volcanic rocks:

(1) The Xiaogulihe ultrapotassic volcanic rocks show relatively ho-mogeneous Os isotopic compositions ranging from 0.1187 to0.1427. The correlation between 187Os/188Os and 1/Os suggeststhat their primary magma was contaminated by negligiblecrustal materials (less than 1%) en route to the surface.

(2) The Xiaogulihe ultrapotassic volcanic rocks mainly originatedfrom phlogopite-bearing garnet peridotite within the lowerSCLM at depths of 80–120 km which had been metasomatizedby potassium-rich silicate melts. The metasomatic potassium-rich silicate melts were mainly generated from the ancientcontinental-derived sediments (N1.5 Ga) that had been subductedand accumulated in the mantle transition zone.

(3) The relatively low 87Sr/86Sr ratios of the Xiaogulihe ultrapotassicrocks imply that their mantle source has evolved with low Rb/Srratios resulting from the metasomatized melts derived from theancient subducted sediments.

Supplementary data to this article can be found online at http://dx.doi.org/10.1016/j.lithos.2014.08.026.

Acknowledgments

Yan Xiao, Pengfei Zhang, Hong Yu and Bin Zhu are sincerely thankedfor their helpful discussions and constructive suggestions. YuehengYang, Yusheng Zhu, Yang Li, Chaofeng Li, Yan Yan, He Li, HongyueWang, Dingshuai Xue, Qian Mao and Yuguang Ma are thanked fortheir help on MC-ICP-MS, TIMS, XRF and EPMA analyses. This workwas financially supported by the National Natural Science Foundationof China (Grants 41173045 and 91214203). Haibo Zou and an anony-mous reviewer are sincerely thanked for their constructive commentswhich have significantly improved the quality of this paper.

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